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Plate convergence usually commences with intra-oceanic Subduction, Andean margins commonly start after ophiolite obduct

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Plate convergence usually commences with intra-oceanic

Subduction, Andean margins commonly start after ophiolite

obduction and subduction flip.

CONVERGENT PLATE MARGINS

Intra-oceanic (ensimatic) subduction

Andean margins

3) Continent - continent collision zones

1)

2)

3)

REMEMBER, IN 3-D A CONVERGENT MARGIN MAY

HAVE DIFFERENT MATURITY ALONG STRIKE!

legend and estimates of plate-tectonic forces

- Fsp - Slab-pull(+)
- Frp - Ridge-push(+)
- Fsu - Suctional force (+)
- For - Orogenic spreading(+)
- Fdf - Mantle drag-force (+ or -)
- Fsr - Subduction ressistance (-)
- Fcd - Extra continental-drag(-)
- Ftr - Transform resistance (-)

Frp = g e (m – w) (L/3 +e/2) ≈ 2*1012 Nm-1

Can also be expressed as a function of age:

Frp = gmTt [1 + (m/(m-w)) 2T/] = 1.19x10-3 t (Unit MPa)

g - gravity ≈ 9,8 ms-2

e – elevation of spreading ridge above cold plate ≈ 3,3 km

(e- is a function of the age [t])

m – mantle density, ≈ 3,2 g cm-3w – water density

L – lithosphere thickness ≈ 85km

T - temperature(~1200C), -thermal diffusivity [ms], - coefficient of thermal expansion [ = 3*10-5 K-1]

Estimate of slab-pull force Fsp pr. unit length subduction zone (see Fowler: Solid Earth, Chap 7, for details)

2z

8gm T1L2Re

2d

Fsp(z) =

[exp(-

) - exp(-

)]

= ca 2x1013Nm-1

4

2ReL

2ReL

- z – depth (d = z give Fsp = 0)
- – coefficient of thermal expansion
- T1 – mantle temp,
- d+L – thickness of the upper mantle
- L – Lithosphere (plate) thickness
- Re – Thermal Reynolds number

Re =(mcpvL)/2k

Thermal Reynolds number

k - conductivity

cp - spesific heat

k - kinematic viscosity

v - subd. velocity

Reactions and phase transitions affecting the forces in subduction zones

- In addition to the thermal contraction and density change will the forces of the subducting litosphere be affected by
- Gabbro to eclogite transition (+)
- Olivin-spinel transiton (+)
- Spinel to oxides (perovskitt and periklas) (-)

Temperature variation across a subduction zone

- Notice the localization of the olivin-spinel and spinel-oxide transitions.
- Use the next fig to explain the phenomena

Phase diagrams for the transititions for olivin to spinel and spinel to post-spinel (oxides)

THE ANATOMY OF A SUBDUCTION COMPLEX

Outer

non-volcanic

island

Fore-arc

basin

Active

volcanic

arc

Remnant-arcs

from

arc-splitting

Back-arc

basin/spreading

alternating

compression

amd tension

Compression

Tension

Trench

sea level

High geotherm

Low geotherm

Accreationary

prism

High geotherm

Please notice that Benioff zones frequently have an irregular shape in 3-D (ex.

Banda Arc). 80% of all seismic energy is released in Benioff zones.

The low geotherm in subductions zones makes them the prime example of high P -

low T regional metamorphic complexes. The high geotherm in the arc-region gives

contemporaneous high-T low P regional metamorphism, together these two regions

give rise to a feature known as”Paired Metamorphic Belts”

Blueshists normally

originate here!

Ophiolites normally

originate here!

alternating

compression

amd tension

Compression

Tension

Trench

sea level

High geotherm

Low geotherm

High geotherm

Seismic quality factor (Q): The ability to transmitt seismic energy

without loosing the energy. Low Q in high-T regions.

Seismic quiet zones---NB potential build-up to very large quakes!

Arc-splitting - tensional regime above subductions zones. Subduction

roll-back.

High heat-flow in the supra-subductions zone regime give rise to

relatively low shallow sealevel above the back-arc basins. Most

ophiolite complexes have their origin is a supra-subduction environment

The amount of energy radiated by an earthquake is a measure of the potential for

damage to man-made structures. Theoretically, its computation requires summing

the energy flux over a broad suite of frequencies generated by an earthquake as

it ruptures a fault. Because of instrumental limitations, most estimates of energy

have historically relied on the empirical relationship developed by Beno Gutenberg

and Charles Richter:

log10E = 11.8 + 1.5MS

where energy, E, is expressed in ergs.

The drawback of this method is that MS is computed from an bandwidth between

approximately 18 to 22s. It is now known that the energy radiated by an

earthquake is concentrated over a different bandwidth and at higher frequencies.

With the worldwide deployment of modern digitally recording seismograph with

broad bandwidth response, computerized methods are now able to make accurate

and explicit estimates of energy on a routine basis for all major earthquakes. A

magnitude based on energy radiated by an earthquake, Me, can now be defined,

Me = 2/3 log10E - 2.9.

For every increase in magnitude by 1 unit, the associated seismic energy increases

by about 32 times.

Although Mw and Me are both magnitudes, they describe different physical properites

of the earthquake. Mw, computed from low-frequency seismic data, is a measure of

the area ruptured by an earthquake. Me, computed from high frequency seismic data,

is a measure of seismic potential for damage. Consequently, Mw [MW = 2/3 log10(MO) - 10.7]

Mw= µ(area)(displacement)]and Me often do not have the same numerical value.

Frictional heating on faults may result in melting of any rock-coposition

Stress-measurements from grain-size and/or dislocation density (4 to 5 x1013m-2)

in olivine associated with pseudotachylytes in peridotite indicate that peridotites

(mantle rocks) may sustain extreme differential stress: 1-3≈ 3-600 MPa.

Assuming a fault with a modest displacement of d ≈ 1m, and a differential

stress of 300 MPa the release of energy according to equation

(1): Wf = Q + E where Q = heat and E = seismic energy is

Wf = d n = d (1-3)/2 =1m(300MPa)/2 ≈ 1.5 x 108 J m-2 or 47 kWhm-2.

The seismic energy (E) is commonly estimated to be < 5% of Wf on a strong fault, ie.

less than 2.3 kWh m-2 is radiated as seismic waves, the remaining energy (Q) turns to

heat and surface energy (difficult to measure) along the fault.

The process is adiabatic since the fault movement occurs in seconds and no heat is lost

by conduction (thermal diffusivity ~1.5 mm2s-1).

Taking the heat capacity of lherzolite, Cp = 1150 J kg-1 oC-1 and a heat of fusion (Fo)

H = 8.6 x106 Jkg-1 the thermal energy (equation 4) required to melt one kg of peridotite:

(4) Q = Cp(T) + H = 1150Jkg-1oC-1 (1200oC) + 8.6 x106 Jkg-1 = 2.7 x 105 Jkg-1.

On a fault with D = 1m, ~60 kg lherzolite may melt pr m2 of the fault plane,

corresponding to an approximately 2 cm thick layer of ultramafic pseudotachylyte.

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